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Clim. Past, 7, 771–800, 2011
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doi:10.5194/cp-7-771-2011
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Climate
of the Past
Deep ocean ventilation, carbon isotopes, marine sedimentation
and the deglacial CO2 rise
T. Tschumi1 , F. Joos1,2 , M. Gehlen3 , and C. Heinze4,5,6
1 Climate
and Environmental Physics, Physics Institute, University of Bern, Sidlerstrasse 5, 3012 Bern, Switzerland
Centre for Climate Change Research, University of Bern, 3012 Bern, Switzerland
3 Laboratoire du Climat et de l’Environnement (LSCE), L’Orme des Merisiers Bât. 712, 91191 Gif-sur-Yvette, France
4 Geophysical Institute, University of Bergen, Allegaten 70, 5007 Bergen, Norway
5 Bjerknes Centre for Climate Research, Bergen, Norway
6 Uni Bjerknes Centre, Uni Research, Bergen, Norway
2 Oeschger
Received: 25 August 2010 – Published in Clim. Past Discuss.: 27 September 2010
Revised: 14 May 2011 – Accepted: 13 June 2011 – Published: 22 July 2011
Abstract. The link between the atmospheric CO2 level and
the ventilation state of the deep ocean is an important building block of the key hypotheses put forth to explain glacialinterglacial CO2 fluctuations. In this study, we systematically
examine the sensitivity of atmospheric CO2 and its carbon
isotope composition to changes in deep ocean ventilation, the
ocean carbon pumps, and sediment formation in a global 3D ocean-sediment carbon cycle model. Our results provide
support for the hypothesis that a break up of Southern Ocean
stratification and invigorated deep ocean ventilation were the
dominant drivers for the early deglacial CO2 rise of ∼35 ppm
between the Last Glacial Maximum and 14.6 ka BP. Another
rise of 10 ppm until the end of the Holocene is attributed
to carbonate compensation responding to the early deglacial
change in ocean circulation. Our reasoning is based on a
multi-proxy analysis which indicates that an acceleration of
deep ocean ventilation during early deglaciation is not only
consistent with recorded atmospheric CO2 but also with the
reconstructed opal sedimentation peak in the Southern Ocean
at around 16 ka BP, the record of atmospheric δ 13 CCO2 , and
the reconstructed changes in the Pacific CaCO3 saturation
horizon.
Correspondence to: T. Tschumi
([email protected])
1
Introduction
Ancient air samples trapped in polar ice reveal that atmospheric CO2 has undergone large natural fluctuations with
amplitudes of 90–100 ppm between glacial and interglacial
periods. Antarctic temperature closely parallels the recorded
atmospheric CO2 concentration during at least the last eight
glacial cycles (e.g., Lüthi et al., 2008).The global oceans
have been identifie
The global oceans have been identified as the key regulator of the atmospheric carbon content on glacial-interglacial
time-scales (e.g., Sigman and Boyle, 2000; Archer et al.,
2000a; Siegenthaler and Wenk, 1984; Sarmiento and Toggweiler, 1984; Knox and McElroy, 1984). The ocean carbon
cycle is driven by a multitude of different processes such as
ocean circulation, the ocean carbon pumps and marine sediment formation which equilibrate with atmospheric CO2 on
time-scales of months to millennia. The main focus of this
study is to examine the millennial-scale response in atmospheric CO2 to changes in the ventilation rate of the deep
ocean considering the carbon cycle feedbacks from implied
modifications in marine biological production and sediment
formation.
We perform a series of illustrative sensitivity experiments
with the coupled Bern3D+C ocean-sediment carbon cycle
model in which the rate of deep ocean ventilation is systematically varied by changing wind forcing. The time-dependent
response for multiple carbon cycle proxies, such as δ 13 C- and
114 C-signatures in the ocean and the atmosphere but also
Published by Copernicus Publications on behalf of the European Geosciences Union.
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T. Tschumi et al.: Deep ocean ventilation, carbon isotopes, marine sedimentation and the deglacial CO2 rise
the depth of the CaCO3 saturation horizon and the sedimentation rate of CaCO3 , organic carbon and opal are compared
to available paleo-data to assess the importance of changes
in deep ocean ventilation for glacial-interglacial CO2 variations.
Paleo-records of CO2 carbon isotope composition provide valuable insight into past sources and sinks of the
atmospheric carbon reservoir. This information can be
used to constrain contributions from the individual processes
made responsible for the CO2 fluctuations (e.g., Elsig et al.,
2009; Joos et al., 2004; Hughen et al., 2004; Smith et al.,
1999; Leuenberger et al., 1992). Unfortunately, during
glacial-interglacial transitions, there were generally several
simultaneous processes generating similar imprints on CO2 ,
δ 13 CCO2 and 114 CCO2 . So far, the isotope records of atmospheric CO2 alone have not allowed for a unique quantitative
process identification (Köhler et al., 2006a).
The timing and coevolution of the CO2 carbon isotope
records over the last deglacation nevertheless indicate that
the initial CO2 increase at the beginning of deglaciations was
caused by the injection of some old carbon stock with low
δ 13 C signature into the atmosphere, as both records of atmospheric δ 13 CCO2 and 114 CCO2 exhibit a relatively rapid
decline during the early deglaciation. Such a carbon release
could in principle come from terrestrial or oceanic sources.
Zeng (2003) proposes that there is a 547 GtC carbon release
from glacial to interglacial originating from organic carbon
under the great ice sheets of the Northern Hemisphere. However, the observed atmospheric CO2 variations are not in
phase with reconstructed sea level fluctuations. Recorded
atmospheric CO2 concentrations show instead an extraordinary correlation with Southern Ocean (SO) climate changes.
It thus appears more plausible that the dominant internal
forcing feedback driving the deglacial CO2 rise is located in
the Southern Ocean (e.g., Fischer et al., 2010).
There is evidence that low glacial CO2 concentrations can
be at least partly attributed to an increased isolation of deep
water masses from the atmosphere (e.g., Skinner et al., 2010;
Boyle, 1988). Many authors postulate that during glacial
times, there existed a relatively large and poorly ventilated
abyssal water mass which led to a relocation of carbon from
the atmosphere to the deep ocean. As most deep ocean watermasses are ventilated through the SO surface (e.g., Primeau,
2005), physical processes in this region are put forth as the
responsible drivers of the postulated changes in deep ocean
overturning (e.g., Fischer et al., 2010; Sigman and Boyle,
2000). It has been proposed that variations in deep ocean
ventilation have resulted from (i) changes in SO wind forcing
(Toggweiler et al., 2006), (ii) changes in SO buoyancy forcing (Watson and Naveira-Garabato, 2006) or (iii) changes
in SO water column stratification (Schmittner and Galbraith,
2008; Francois et al., 1997). However, there is no broadly accepted consensus so far on the individual role of the different
physical mechanisms nor on the magnitude of the associated
CO2 variations during glacial-interglacial transitions.
Clim. Past, 7, 771–800, 2011
Reconstructions of the glacial δ 13 C distribution from benthic foraminifera in the Pacific Ocean show that – relative
to the Holocene – deep water below 2000–2500 m was more
depleted in δ 13 C compared to the upper water mass (Matsumoto et al., 2002; Keigwin, 1998). In the Indian, Atlantic
and Southern Oceans, reconstructed glacial-age vertical δ 13 C
gradients point at the existence of a sharp chemocline at similar depths separating the well-ventilated water above from
poorly ventilated water below (e.g., Hodell et al., 2003; McCorkle et al., 1998; Curry et al., 1988; Duplessy et al., 1988).
These δ 13 C-data support the “nutrient deepening” hypothesis (Boyle, 1988) calling for a transfer of nutrients and DIC
from upper- and mid-depth waters to the deep ocean driven
by a more efficient marine biological pump during glacial
times. Several potential triggers for an increased biological pump efficiency during glacials have been suggested in
the literature. For instance, reduced deep-water ventilation
rates (e.g., Hodell et al., 2003), slowed organic matter degradation under colder water temperatures (Matsumoto et al.,
2007; Kwon et al., 2009) or significantly increased aeolian
iron supply to the ocean surface (Martin, 1990).
Jaccard et al. (2009) and Bradtmiller et al. (2010) have
recently presented evidence for increased sequestration of
remineralized carbon into the deep Pacific during the LGM.
Their reasoning is based on the argument that a more efficient
biological pump during glacial periods does not necessarily
imply that total (= preformed + regenerated) nutrient concentrations in the deep ocean were significantly higher than during interglacials. Rather, the increase in biological pump efficiency would be mirrored in lower oxygen concentrations in
the deep ocean together with an increase in regenerated nutrients which might be at least partially offset by a decrease
in the preformed nutrient component (e.g., Ito and Follows,
2005).
Further evidence for a slowed deep ocean ventilation
during glacial periods comes from recorded variations of
opal burial in SO sediments (Anderson et al., 2009) and
the occurrence of δ 13 C carbon isotope minima in planktic
foraminifera at the beginning of glacial terminations (Spero
and Lea, 2002). These signals have been interpreted to reflect an invigoration in SO ventilation during the last glacialinterglacial transition. As a consequence, the breakdown of
surface water stratification and renewed Circumpolar Deep
Water (CDW) upwelling in the SO would bring the postulated abyssal watermass – which was isolated from the atmosphere during the glacial period – to the ocean surface.
The large recorded drop in atmospheric 114 C of about
300 ‰ over the course of the last deglaciation may further
support this interpretation as the drop cannot be attributed
solely to a decrease in stratospheric 14 C-production (Hughen
et al., 2004; Muscheler et al., 2004). Radiocarbon measurements on glacial-age sediments from the North Pacific (Galbraith et al., 2007; Marchitto et al., 2007) and from the deep
Atlantic (Keigwin and Schlegel, 2002; Boyle, 1988) indicate that abyssal waters were indeed less ventilated during
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T. Tschumi et al.: Deep ocean ventilation, carbon isotopes, marine sedimentation and the deglacial CO2 rise
glacial times. However, surface-to-deep 14 C-age differences
reconstructed from a sediment core in the western equatorial Pacific show that late-glacial ventilation at a depth
of 2.8 km was not significantly lower than today (Broecker
et al., 2007). Summarizing the available evidence from carbon isotopes, Broecker and Barker (2007) come to the conclusion that the volume of an isolated abyssal reservoir during glacials is limited to no more than half of the ocean.
The “shelf hypothesis” (Broecker, 1982) is an alternative
oceanic mechanism that has been put forth to explain the
recorded glacial-interglacial CO2 fluctuations. This mechanism relies on changes in the biological pump turnover
driven by shifts in organic matter accumulation on shallow
shelves as sea level rises and falls. Atmospheric pCO2 would
be reduced during glacial times through increased biological
productivity promoted by an enhanced whole-ocean nutrient
inventory. Paleo-data have revealed, however, that several
prerequisites are not met which are required for the shelf hypothesis to explain the full glacial CO2 drawdown: For instance, there is no sufficiently large amount of organic sediments on shelf areas to provide additional nutrients during
glacials (Peacock et al., 2006). Recent studies have suggested that an increase in the glacial nutrient inventory might
be linked to changes in ventilation at intermediate depths
rather than to changes in sea-level. A decreased volumetric extent of oxygen minimum zones would have reduced
both the marine sinks for NO−
3 (Deutsch et al., 2004) and
2010).
for PO3−
(Wallmann,
4
Besides reorganizations of the ocean-internal carbon cycle and the terrestrial biosphere, transient imbalances between marine sediment formation and continental weathering fluxes may also have contributed significantly to the
recorded millennial-scale variations in atmospheric CO2 ,
δ 13 CCO2 and 114 CCO2 (e.g., Archer et al., 2000a). Most
prominently, the mechanism of carbonate compensation may
have accounted for up to a third of the glacial-interglacial
CO2 change without conflicting with reconstructed seafloor
CaCO3 preservation patterns (Marchitto et al., 2005). Imbalances in the marine organic matter budget have received less
attention in previous studies. Nevertheless, there is a priori
no reason why excess burial or excess weathering of organic
matter should not also have affected atmospheric CO2 levels
in the past.
Numerical models allow a process-based interpretation of
the recorded CO2 variations and to quantitatively test different scenarios of glacial-interglacial CO2 variations. A
dilemma is that the application of comprehensive climatecarbon models for multi-millennial simulations and for exploring the sensitivity of results to uncertain parameters is restricted by their large CPU-demand. On the other hand, box
models are inherently limited as they do not include ocean
dynamics and only poorly represent the 3-D geometry of the
ocean and ocean-sediment interactions.
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The goal of this study is to explore the carbon cycle implications of altered deep ocean ventilation and biological
carbon pumps together with the millennial-scale effects of
implied changes in marine sediment formation. For this purpose, we use the global coarse-resolution 3-D Bern3D+C
ocean model (Müller et al., 2006; Tschumi et al., 2008),
which is coupled to a spatially-resolved dynamic model for
marine sediment formation (Tschumi, 2009). This modelling
framework allows for the simulation of multiple tracers over
long time periods of several 10 000 years which are required
for sediment formation processes to relax to equilibrium.
The model simulates the cycling of total carbon
(12 C + 13 C + 14 C) and the carbon isotopes 13 C and 14 C between the atmosphere, the terrestrial biosphere, the ocean
reservoir, and the marine sediments. The ocean circulation field, the ocean carbon pumps and the formation of organic, opal, and CaCO3 sediments at the seafloor are calculated from prognostic dynamical equations discretized on
3-D grids. The terrestrial reservoir and the atmosphere are
represented rather simplistically as single well-mixed boxes.
The carbon stock on land is fixed (2220 GtC), but not its isotopic composition.
As a starting point of this study, we have calculated the
response of the coupled model to a change in the export ratio of organic matter to calcite (Sigman et al., 1998). This
experiment served as a test for the model representation of
carbonate compensation and allow us to compare our modelling framework to others. The second model experiment
was an increase of the whole-ocean nutrient inventory by
30 % (Sigman et al., 1998; Broecker and Peng, 1987) serving
as a test for the “shelf hypothesis”. Note that a 30 %-increase
in whole-ocean nutrients during glacials is likely to be an
upper limit of a realistic change (Deutsch et al., 2004; Wallmann, 2010). In contrast to previous similar experiments,
in our model stimulated export production of organic matter
will eventually relaxed back to initial rates as excess organic
matter sedimentation removes the injected nutrients from the
ocean in the long term. Therefore, the modeled shift in the
Pacific CaCO3 saturation horizon in our experiment turned
out to be small as opposed to previous studies (Heinze et al.,
1991; Sigman et al., 1998). Finally, but as the main interest
of this study, we have probed the response to variations in
ocean deepwater ages driven by changes in SO wind forcing
and the implied changes in SO deep convection for a range
of different ocean ventilation time-scales. All of these sensitivity experiments started from the same model steady-state
that is compatible with preindustrial boundary conditions for
ocean physics and the biogeochemical cycles as described in
Sect. 1.2.
The novel aspect of our analysis is the simulation of the
time-dependent response to prototype scenarios considering
13 C- and 14 C-cycling explicitely in a fully-coupled dynamical ocean-sediment model with sedimentation of all particulate species, i.e. the marine organic matter, CaCO3 and opal
cycles are all represented as “open cycles” in this study. The
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broad range of processes and variables simulated increases
the complexity of the ocean carbon cycle in the model, but
it also allows for a more robust simultaneous comparison
of the model results with multiple proxies. The systematic
variations in deep ocean water ages through changes in the
wind forcing and deep convection in the Southern Ocean represents a new approach to explore the sensitivity of atmospheric CO2 to deep ocean ventilation and marine sediment
formation in 3-D models.
1.1
Model description
The version of the Bern3D+C carbon cycle model used here
comprises an atmospheric, a terrestrial, an oceanic and a sedimentary reservoir which are coupled through the exchange
of carbon and its isotopes 13 C and 14 C. The model’s ocean
physics is based on the work of Edwards et al. (1998) as updated by Edwards and Marsh (2005) and Müller et al. (2006).
The marine biogeochemical cycles are described by Parekh
et al. (2008) and Tschumi et al. (2008) and the coupling to
the sedimentary component in Tschumi (2009). Unless explicitly stated otherwise, the model is run under standard parameters as mentioned in these references.
Marine biogeochemical cycles and air-sea gas-exchange
are implemented following OCMIP-2 (Orr et al., 1999; Najjar et al., 1999) with the addition of prognostic formulations
for marine biological productivity as well as representations
for the cycling of iron (Parekh et al., 2008), silica (Tschumi
et al., 2008), 13 C and 14 C.
Export production is a function of temperature, light,
phosphate and iron following Doney et al. (1996). The
competition between opal and calcite producers is modeled
following formulations of the HAMOCC5 model of MaierReimer et al. (2005). The half-saturation constant for silicauptake is a function of local silica concentration and the ratio
of silica to phosphorous uptake by opal producers varies with
silica, light, iron and phosphate following Aumont and Bopp
(2006).
The atmosphere is considered as a single well-mixed box.
To calculate air-sea gas-exchange, a constant value for the
square of the wind-speed u2 is employed which is set equal to
the spatio-temporal mean of the climatology for u2 provided
by OCMIP-2 (Najjar et al., 1999).
The land biosphere model taken from Siegenthaler and
Oeschger (1987) is composed of 4 well-mixed compartments with fixed carbon exchange fluxes and stocks: ground
vegetation and leaves (100 GtC), wood (500 GtC), detritus
(120 GtC) and soils (1500 GtC).
The sedimentary component is based on the model described by Gehlen et al. (2006) and Heinze et al. (1999)
which represents sediment diagenesis in the top 10 cm beneath seafloor. Sediment formation, i.e. the accumulation
of opal, CaCO3 , organic matter and clay sediments is calculated on the basis of a set of dynamical equations for
the sedimentary diagenetic processes. The redissolution and
Clim. Past, 7, 771–800, 2011
remineralization reactions all follow first order kinetics. Two
types of remineralization reactions are considered, oxic respiration and denitrification. For more details on the sediment
model component and model equations, please refer to the
appendix.
In the model, cycling of 13 C is subject to the following
fractionation effects: in the ocean, equilibrium fractionation is considered in the chemical carbonate equilibrium system (Mook, 1988). For air-sea gas-exchange, a kinetic fractionation is computed (Siegenthaler and Muennich, 1981).
Further fractionations occur during marine photosynthesis
(Freeman and Hayes, 1992) and the formation of calcium
carbonate (Mook, 1988). In the land biosphere, isotopic fractionation is accounted for during photosynthesis.
Cycling of 14 C is simulated explicitly, i.e. the actual concentrations of 14 C are transported and not the 14 C/12 C-ratio
(Orr et al., 1999).
1.2
Model spin-up and preindustrial steady state
The coupled carbon cycle model was spun up by bringing the
terrestrial biosphere, the ocean and the sediments into steady
state under fixed atmospheric values for pCO2 (278.00 ppm),
δ 13 C (−6.305 ‰) and 114 C (0.00 ‰). A fixed production
rate for 14 C was then prescribed and pCO2 , δ 13 C and 114 C
were let evolve dynamically in the atmosphere to perform
the subsequent sensitivity simulations. The remaining drifts
in the control simulation was less than 3 × 10−3 ppm kyr−1
in pCO2 , less than 2 × 10−5 ‰ kyr−1 in δ 13 C and less than
3 × 10−3 ‰ kyr−1 in 114 C. As the typical model response
to the perturbations applied was much larger than that, the
spun up model state is sufficiently close to equilibrium for
our purposes.
The combined ocean/sediment compartment was separately intialized as follows before it was coupled interactively to the atmosphere. To spin up ocean physics in a
first step, sea surface temperatures (SST) and salinities (SSS)
were restored towards modern monthly observations (Levitus et al., 1994; Levitus and Boyer, 1994) and wind stress
was prescribed according to the monthly climatology from
NCEP/NCAR (Kalnay and et al., 1996). After 10 000 model
years, mixed boundary conditions were imposed for SST and
SSS following Tschumi et al. (2008). Inspection of the modeled ocean circulation field and the temperature and salinity
distributions at steady state revealed that the primary flow
paths of the global ocean circulation and the important water masses were represented adequately (see Müller et al.,
2006; Tschumi et al., 2008). The annual mean maximum Atlantic meridional overturning circulation (AMOC) amounted
to 14.7 Sv, comparing favorably to the estimate of 17 ± 5 Sv
(Talley et al., 2003).
In a second step, the steady state ocean circulation was
then used to spin up marine biogeochemical cycles during
10 000 model years. The resulting steady state fields of the
modeled biogeochemical tracers successfully reproduced the
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T. Tschumi et al.: Deep ocean ventilation, carbon isotopes, marine sedimentation and the deglacial CO2 rise
large-scale patterns of the observed distributions in the modern ocean (see Tschumi et al., 2008; Müller et al., 2006,
and Fig. 3). Simulated export fluxes for particulate organic
matter (12.76 GtC yr−1 ), CaCO3 (1.13 GtC yr−1 ) and opal
(95 Tmol Si yr−1 ) all lay within available estimates (Jin et al.,
2006; Sarmiento and Gruber, 2006).
In a next phase, the sediments were “filled up” during
50 000 years by imposing spatially uniform, but temporally
variable weathering fluxes to the sea surface for phosphate,
silicic acid, alkalinity, carbon, 13 C and 14 C. The weathering rates were set equal to the net flux of the respective
tracers across the ocean-sediment interface as diagnosed at
each time-step such that annual mean ocean inventories are
kept practically unchanged. As the sediment inventories converged to their steady state values, weathering rates stabilized
to balance with the loss of solid matter that results from sediment burial. These values were diagnosed and prescribed
subsequently as temporally constant weathering input. The
ocean-sediment component was run into a new equilibrium
during 30 000 years in this new setup.
Simulated sediment accumulation amounted to
0.14 GtC yr−1 for organic matter and to 0.17 GtC yr−1
for CaCO3 . These rates are roughly compatible with
data-based estimates of 0.12–0.26 GtC yr−1 for organic
matter burial and of 0.1–0.14 GtC yr−1 for CaCO3 burial
(Sarmiento and Gruber, 2006; Feely et al., 2004). In a series
of carbon pulse release experiments, CO2 neutralization
time scales were found to be between 5–8 kyr for seafloor
CaCO3 neutralization and between 8–13 kyr for terrestrial
CaCO3 neutralization (Tschumi, 2009). These values are
compatible with previous studies (e.g., Archer et al., 1997).
The model equations, the model parameters and the details
of the ocean-sediment coupling can be found in Tschumi
(2009).
Simulated sediment composition in the spun up coupled
steady state is shown in Fig. 1a–c. Table 1 displays bulk
numbers of the simulated particulate fluxes exported from the
sea surface, deposited at the sea floor and accumulated in the
sediments along with available estimates for the preindustrial
time
2
2.1
Model experiments and results
Rain ratio reduction
The marine biological pumps (Volk and Hoffert, 1985),
i.e. the organic matter and the carbonate pumps, are central drivers of the ocean carbon cycle. For two separate
reasons, they are of great importance for the regulation of
the atmospheric CO2 level by the oceans. First, they establish systematic gradients in dissolved inorganic carbon
(DIC) and alkalinity between the surface and the deep ocean.
The overall impact of both pumps together is lower pCO2
levels in the surface ocean compared to an abiotic ocean.
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Second, accumulation of organic matter and CaCO3 sediments is driven by the deposition of biogenic material on the
seafloor. Any change in the marine biological pumps thus
also affects organic matter and CaCO3 sedimentation with
potentially significant impacts on atmospheric CO2 through
perturbations of the marine organic matter and CaCO3 balances.
The ocean carbonate pump is driven by the production
of biogenic CaCO3 shells in the sea surface. The sinking
CaCO3 shells entrain alkalinity to the deep ocean and hence
increase sea surface pCO2 by shifting the speciation of the
surface ocean DIC towards dissolved CO2 . A reduced export ratio between CaCO3 and organic matter export (rain
ratio) would reduce sea surface pCO2 , not only due to an
ocean-internal redistribution of alkalinity but also as a result
of a lower carbonate concentration in the deep ocean triggering redissolution of marine CaCO3 sediments. The resulting
imbalance between the weathering flux of dissolved CaCO3
into the ocean and marine CaCO3 sedimentation would drive
a rise in the ocean’s total alkalinity and additionally contribute to lower sea surface pCO2 . Carbonate compensation
is the feedback mechanism responsible for re-establishing
the ocean CaCO3 balance on a time-scale of order 10 kyr
through shifts in the depth of the CaCO3 saturation horizon
(Archer et al., 2000a).
By reason of their central role in the regulation of the
ocean’s alkalinity budget, rain ratio mechanisms were among
the first candidates made responsible for driving glacialinterglacial CO2 fluctuations in the atmosphere (e.g., Archer
et al., 2000a; Berger and Keir, 1984). A reduced rain ratio
during glacial times may have resulted, for instance, from increased diatom productivity in lower latitudes at the expense
of coccolithophorids (silica leakage hypothesis, Matsumoto
et al., 2002) or from relaxed nutrient-limitation in the sea surface (Margalef, 1978). The inferred CaCO3 preservation history recorded in marine sediments, however, constrains the
contribution of whole-ocean alkalinity changes to roughly a
third of glacial-interglacial CO2 variations (Marchitto et al.,
2005).
In our first model experiment, we attempt to disentangle
and quantify the complex response in ocean biogeochemistry
to a reduced rain ratio. We perform parallel simulations in a
coupled ocean-sediment setup and in an ocean-only setup to
separate ocean-internal mechanisms from interactions with
the sediments. Following Sigman et al. (1998), we abruptly
reduce CaCO3 export by 20 % uniformly in the global sea
surface. This corresponds in our model to a decline in the
global rain ratio from 0.088 to 0.071.
2.1.1
Results
Figure 3 shows the time-dependent carbon cycle responses
as simulated with the ocean-only model (black lines) and the
ocean-sediment model (red lines). In response to decreased
CaCO3 , export atmospheric CO2 drops from 278 ppm to
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Fig. 1. Modelled (left row) and observed (right row) composition of surface sediments on the ocean bottom: (a) particulate organic matter,
(b) CaCO3 and (c) opal. Modelled sediment composition is averaged over the top 1.1 cm of the sediment column. Sediment data is from
the Pangaea database (publishing Network for Geoscientific & Environmental Data, http://orfois.pangaea.de). Note that the color scale is not
linear.
268 ppm within centuries (e-folding time scale: 390 yr) in
the ocean-only model (Fig. 3a, black line). Mean surface alkalinity rises as less CaCO3 shells sink from the surface to
the ocean interior (Fig. 3b, black line). Reduced alkalinity in
the ocean interior in turn lowers [CO2−
3 ] pushing the CaCO3
saturation horizon upwards by approximately 300 m (Fig. 3d,
black line).
Clim. Past, 7, 771–800, 2011
In the ocean-sediment model, the CO2 drawdown is
strongly amplified to 237 ppm due to the increase in wholeocean alkalinity (Fig. 3a and b, red line). Carbonate compensation responds rather slowly with an e-folding time scale
of 11.7 kyr. Reduced CaCO3 export and deposition on the
seafloor drive a deepening in the CaCO3 saturation horizon
as to provide a larger area on the seafloor where the rain of
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777
Fig. 2. Modelled distribution of the isotope signatures in ocean dissolved inorganic carbon for preindustrial boundary conditions. 114 CDIC
is shown in the left row and δ 13 CDIC in the right row. Top panels (a) and (d) show bottom water signatures, the middle row panels (b) and (e)
display Atlantic and bottom panels (c) and (f) Pacific zonal mean sections in preindustrial steady state.
CaCO3 is efficiently preserved (Fig. 3d, red line). As a consequence, the simulated CaCO3 sedimentation rate starts to
recover after the initial abrupt decline in order to restore the
ocean’s CaCO3 balance in the model (Fig. 3c). The final total shift in steady state saturation depth in the ocean-sediment
model amounts to a deepening of 700 m in the Atlantic and
of 800 m in the Pacific basin (Fig. 3d, red line).
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The simulated responses in atmospheric δ 13 CCO2 and
to changes in the rain ratio are small. The
ocean-internal distribution of δ 13 CDIC and 114 CDIC is only
marginally affected by the marine CaCO3 cycle. In the
ocean-only experiment, we therefore simulate negligible
changes in atmospheric δ 13 CCO2 and 114 CCO2 as CaCO3
export is reduced. In fact, there is virtually no response in
114 CCO2
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T. Tschumi et al.: Deep ocean ventilation, carbon isotopes, marine sedimentation and the deglacial CO2 rise
Fig. 3. Rain ratio reduction: Simulated response to a 20%-reduction in CaCO3 export. (a) Atmospheric CO2 , (b) mean surface alkalinity,
(c) CaCO3 sedimentation, (d) shift in Pacific CaCO3 saturation horizon, (e) atmospheric δ 13 CCO2 and (f) atmospheric 114 CCO2 .
atmospheric δ 13 CCO2 (Fig. 3e, black line) and 114 CCO2 is
raised by a modest 3 ‰ (Fig. 3f, black line).
In the coupled ocean-sediment experiment, the system response in carbon isotope signatures of atmospheric CO2 is
additionally sensitive to the decline in marine CaCO3 sedimentation. However, the responses of δ 13 CCO2 and 114 CCO2
to carbonate compensation are rather slow and remain small
(Fig. 3e and f, red lines).
2.1.2
Comparison to other studies
Heinze et al. (1991) have simulated a relatively modest sensitivity of atmospheric CO2 to changes in the rain ratio. They
have found a CO2 drawdown of 28.5 ppm when the rain ratio is reduced by 50 % in the 3-D Hamburg Carbon Cycle
Model which was coupled to a simplistic model for sediment
diagenesis with fixed sediment resuspension rates. Archer
and Maier-Reimer (1994) have shown that including organiccarbon-driven dissolution of CaCO3 sediments substantially
Clim. Past, 7, 771–800, 2011
increases the pCO2 response to rain ratio changes in their
model. They reach glacial CO2 levels by reducing the rain
ratio by 40 %, which amounts to a CO2 sensitivity comparable to this study.
Sigman and Boyle (2000) have reduced the ratio of organic
matter to CaCO3 export in the low latitudes by 50 % in the
CYCLOPS box model. They obtain a drawdown of 22 ppm
in the ocean-only model and of 63 ppm in the ocean-sediment
model. CaCO3 compensation amplifies the CO2 reduction by
a factor of three in their study compared to a factor of four
in this study. The reason for the comparatively lower impact
of CaCO3 compensation in the box model study is probably
the shorter relaxation time of CaCO3 burial. The original
rate is restored after about 15 kyr, whereas in the Bern3D+C
model it takes roughly 50 kyr. The longer relaxation time
in our study results in a larger imbalance between CaCO3
burial and weathering when integrated over time such that
the increase in whole-ocean alkalinity is larger in response to
carbonate compensation.
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Chikamoto et al. (2009) have lowered the rain ratio by
25 % in the High-Latitude Exchange/Interior Diffusion Advection (HILDA) model coupled to a similar dynamical sediment model as used in this study. They find a drawdown
of 59 ppm in atmospheric CO2 together with a deepening of
the calcite saturation horizon by 1870 m. The response time
scale of atmospheric CO2 and the magnitude of the drawdown are comparable to the results of our study; however,
we find a deepening of the order of only 800 m in the CaCO3
saturation horizon.
In conclusion, it appears to be challenging to robustly
quantify the effect of rain ratio changes on atmospheric CO2
by means of carbon cycle models. The simulated CO2 response depends on the specific biogeochemical model state,
such as the initial rain ratio or the CaCO3 throughput in the
oceans, as well as on the model parametrization of the oceansediment interactions. Our results are consistent with current understanding, suggesting that carbonate compensation
would amplify the CO2 response to changes in the ocean carbonate pump by a factor of as much as 3–4.
2.2
Strengthening of the marine biological cycle
As an alternative to a reduced rain ratio, low atmospheric
CO2 during glacial periods might have resulted from more
vigorous biological carbon cycling in the ocean. A strengthened marine biological cycle might be driven for instance
by increased nutrient availability (Broecker and Peng, 1982),
iron fertilization (Martin and Fitzwater, 1988) or changes in
nitrogen cycling (Falkowski, 1997).
In the next experiment, we examine the response in atmospheric CO2 , δ 13 CCO2 and 114 CCO2 to a stimulated ocean
biology by increasing the marine nutrient inventory in the
Bern3D+C model. The experiment is again performed with
and without coupling to sediments to isolate the effect of
ocean-sediment interactions.
Following the procedure of Archer et al. (2000a); Sigman
et al. (1998) and Broecker and Peng (1987), we have abruptly
increased PO4 concentrations uniformly by 30 % to stimulate simultaneously both the organic matter and the carbonate pumps. With our model biology the global rain ratio remained virtually unchanged. Therefore, the sedimentary response on atmospheric CO2 is dominated in our model by
excess organic matter sedimentation rather than by carbonate compensation.
2.2.1
Atmospheric CO2
In the ocean-only model, the augmented PO4 inventory
drives a persistent increase in organic matter and CaCO3 export production both by roughly 11 % (Fig. 4c and d, black
lines). The elevated rates of organic matter and CaCO3 export have opposite effects on sea surface pCO2 . As the
strengthened organic matter pump dominates the net ocean
response, atmospheric CO2 gets reduced to 255 ppm with an
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e-folding time scale of 470 years in the ocean-only experiment (Fig. 4a, black line). As the concurrent changes in
downward fluxes of POC and CaCO3 have roughly opposing effects on [CO2−
3 ] at the depth of the saturation horizon,
there is only a small shift in calcite saturation depth of less
than 20 m in the ocean-only model (Fig. 4b, black line).
Biological export of carbon from the sea surface augments
primarily in the Atlantic and in the Indian oceans, and to
a lesser degree in the Northern Pacific. There is almost no
change in export production in the equatorial and South Pacific as well as in the Southern Ocean (south of 51◦ S) because biological productivity in these regions is limited by
iron or light availability. Note that a glacial increase in export production in the North Pacific is at odds with paleoproductivity reconstructions (see Jaccard et al., 2009; Galbraith
et al., 2007; Crusius et al., 2004), suggesting that the history
of export production from the North Pacific surface was predominantely controlled by processes other than changes in
the whole-ocean nutrient inventory such as changes in water
column stratification and associated upwelling of nutrients
from the deep (Galbraith et al., 2007).
In the coupled ocean-sediment setup, atmospheric CO2 is
reduced to 232 ppm within 30 kyr (Fig. 4a, red line). Thereafter, CO2 slightly increases by 2 ppm until the full balance
between weathering and burial of organic matter and CaCO3
is re-established after 50 ky (Fig. 4e and f). The effect of
ocean-sediment interactions thus enhances the CO2 drawdown in the coupled system by a factor of two.
The nutrient increase abruptly boosts the export of organic
matter and CaCO3 by 8.5 % in the ocean-sediment model
(Fig. 4c and d, red lines). Subsequently, both export rates decline and reach almost original values after 50 kyr. As both
changes in CaCO3 and organic matter burial are taken into
consideration, elevated organic matter export is not sustained
in the long run here, in contrast to previous studies (Sigman
et al., 1998; Broecker and Peng, 1987) which have accounted
solely for changes in CaCO3 sedimentation.
Excess organic matter burial draws nutrients out of the
ocean to bring organic sedimentation back to balance the
riverine input of nutrients. The initial increase in the nutrient
inventory is thereby neutralized in the long run. The simulated decline in CaCO3 export also follows directly from
the gradual decrease in nutrient availability. In our model
biology, CaCO3 export is computed from primary productivity and is not affected by changes in the calcite saturation state. The effect of CaCO3 compensation on restoring
CaCO3 burial is therefore only of secondary importance, in
contrast to the rain ratio reduction experiment.
The modeled CaCO3 saturation horizon deepens by about
150 m in the Atlantic and by 300 m in the Pacific during the
first 15 000 years with a subsequent shoaling in both basins
thereafter (Fig. 4b, red line). The first part of the response
is dominated by excess POC burial, the second part by excess CaCO3 burial. Since excess POC burial implies a loss
of whole-ocean DIC, it tends to push the saturation horizon
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Fig. 4. Strengthening of the marine biological cycle: Simulated response to an abrupt 30 %-increase in the ocean’s phosphate content.
(a) Atmospheric CO2 , (b) shift in Pacific CaCO3 saturation horizon, (c) POC export production, (d) CaCO3 export production, (e) POC
sedimentation, (f) CaCO3 sedimentation, (g) atmospheric δ 13 CCO2 and (h) atmospheric 114 CCO2 .
downwards. On the contrary, excess CaCO3 burial leads to a
shoaling of the saturation horizon as alkalinity and DIC are
removed from the ocean in a 2:1-ratio. The final shift in the
calcite saturation horizon after 50 kyr is a deepening of less
than 100 m.
Clim. Past, 7, 771–800, 2011
2.2.2
Atmospheric δ 13 CCO2 and 114 CCO2
The simulated changes in atmospheric δ 13 CCO2 and
114 CCO2 are predominantly controlled by the response in organic matter cycling. In the ocean-only model, atmospheric
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T. Tschumi et al.: Deep ocean ventilation, carbon isotopes, marine sedimentation and the deglacial CO2 rise
δ 13 CCO2 rises by 0.16 ‰ and 114 CCO2 by 7 in response
to the strengthened marine biological cycle (Fig. 4g and h,
black lines). δ 13 CCO2 rises as surface ocean δ 13 CDIC is
shifted towards more positive values due to enhanced production of organic matter. The increase in 114 CCO2 is the
direct consequence of reduced atmospheric CO2 under a constant 14 C-production rate.
In the ocean-sediment model, δ 13 CCO2 increases by
∼0.4 ‰ and 114 CCO2 rises by roughly 70 ‰ in response to
a strengthened marine biological cycle (red lines in Fig. 4g
and h). The time-dependent response in the carbon isotope
signatures of atmospheric CO2 is more complex than in the
ocean-only system. In particular, relaxation times are much
longer due to the slow ocean-sediment interactions.
Next to the (transient) changes in the individual pumps’
strengths, elevated sedimentation rates of organic matter and
CaCO3 additionally affect the response in isotopic carbon
signatures in the ocean-sediment setup. The sedimentary response in atmospheric δ 13 CCO2 is rather involved since excess organic matter burial removes carbon with negative δ 13 C
and excess CaCO3 burial draws carbon with positive δ 13 C
out of the system. Excess organic matter sedimentation dominates the response in atmospheric δ 13 CCO2 during the initial
25 kyr, whereas excess CaCO3 burial drives a slow reduction in δ 13 CCO2 thereafter (red line in Fig. 4g). This timedependent behavior results from the longer relaxation time
of the marine CaCO3 cycle to the intensified marine biology
as compared to the relaxation time of the organic matter cycle.
The sedimentary response pushes atmospheric 114 CCO2
signatures towards more positive values (red line in Fig. 4h).
On the one hand, the ocean-sediment interactions draw CO2
out of the atmosphere which elevates 114 CCO2 under constant 14 C-production. On the other hand, enhanced organic
matter and CaCO3 sedimentation further contribute to rising atmospheric 114 CCO2 as excess burial removes relatively old carbon (with low 114 C) from the combined oceanatmosphere system.
2.2.3
Comparison to other studies
Heinze et al. (1991) have increased the marine PO4 inventory
by 30 % in the 3-D Hamburg ocean-sediment model. They
have found that atmospheric CO2 decreased by 60 ppm and
the Pacific CaCO3 saturation horizon deepened by roughly
900 m. Sigman et al. (1998) only let low-latitude production
rise upon a 30 %-increase in whole-ocean PO4 inventory in
the CYCLOPS box model. They find an increase of 32 %
in low-latitude export production together with a CO2 reduction of 32 ppm in the ocean-only setting. When CaCO3 sedimentation is additionally taken into account, the simulated
CO2 decline is only of 19 ppm. As opposed to our results,
the sedimentary response in their simulation is to counteract the CO2 drawdown resulting from strengthened biological pumps. Sigman et al. (1998) find a mean global shoaling
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in CaCO3 saturation depth of 500 m, whereas here we obtain
only a modest deepening.
The reason for this discrepancy is mainly the fact that POC
burial is not taken into account explicitly by Sigman et al.
(1998). The nutrient-feedback is thus not considered. In our
simulations, the loss of nutrients through excess POC burial
on the one hand draws carbon out of the ocean and on the
other hand limits excess CaCO3 export in such a way that
CaCO3 compensation has only a minor effect opposing the
CO2 reduction. The impact of CaCO3 compensation on atmospheric CO2 is seen in a small increase occurring after
30 kyr (Fig. 4a, red line).
The following conclusions emerge. Perturbations in global
burial fluxes due to ocean processes are removed on a multimillennial time scale to re-establish the balance between removal of organic matter, CaCO3 , and opal by burial and corresponding weathering inputs from rivers and winds. As a
consequence of this sedimentary compensation, some proxy
signals that arise from changes in the internal ocean carbon
cycle are partly or completely muted with time (see also next
Sect. 3.3). Differences in proxies between distant time periods, for instance between the LGM and the late Holocene,
may therefore partly reflect changes in the balance of the
weathering-sedimentation cycle.
Kohfeld et al. (2005) compared productivity records for
the late Holocene, the Last Glacial Maximum and marine
isotope stage 5a to 5d. They concluded that export production was globally higher during the LGM than during either
stage 5a-5d or the Late Holocene. They suggest a limited role
for the marine biological cycle to explain glacial-interglacial
CO2 variations. Our results indicate that such a conclusion
should be viewed with caution, because differences in proxies might also be related to imbalances in the weatheringsedimentation cycle. Potential secular trends arising from
these imbalances should also be considered when analyzing
whole ocean changes in δ 13 C to infer variations in organic
matter pools (Shackleton, 1977; Bird et al., 1994) or in atmospheric δ 13 CCO2 (Leuenberger et al., 1992), as well as when
interpreting the recorded evolution of nitrogen and silicate
isotope signatures (Deutsch et al., 2004; De La Rocha, 2006).
2.3
Global ocean ventilation
The last set of simulations focusses on the response in atmospheric CO2 , δ 13 CCO2 and 114 CCO2 to changes in the rate
of deep ocean ventilation. In these idealized experiments we
vary deep ocean ventilation in the model by uniformly scaling the prescribed amplitude of both zonal and meridional
wind stress in the Southern Ocean (south of 51◦ S). The timedependent model response in atmospheric CO2 , δ 13 CCO2 and
114 CCO2 is then calculated as the result of the interplay between changes in ocean circulation, marine biogeochemical
cycling and sediment diagenesis subsequent to abrupt change
in wind boundary conditions. Note that gas-exchange velocities are kept unaffected and feedbacks mediated by SST
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Table 1. Bulk numbers characterising the preindustrial steady state of the coupled ocean-sediment Bern3D+C model compared to previous
modelling results and independent estimates.
Variable
Units
this study
Heinze et al. (2003)
Ind. estimates
atm. CO2
POC export production
POC rain onto sediments
POC accumulation
CaCO3 export production
CaCO3 rain onto sediments
CaCO3 accumulation
Opal export production
Opal rain onto sediments
Opal accumulation
POC sediment pool
CaCO3 sediment pool
Opal sediment pool
ppm
GtC yr−1
GtC yr−1
GtC yr−1
GtC yr−1
GtC yr−1
GtC yr−1
Tmol Si yr−1
Tmol Si yr−1
Tmol Si yr−1
GtC
GtC
Tmol Si
278.0
12.76
0.49
0.14
1.13
0.40
0.17
95
64.9
1.12
428
881
9779
277.7
8.66
0.12
0.048
1.64
0.42
0.18
185
81.3
5.5
241.9
1010
19 986
–
6.5–13.1a
1.7–3.3a
0.12–0.26a
0.8–1.2b
0.5c
0.1–0.14d
102–178b
13–47e
5.3–8.9e
–
–
–
a Sarmiento and Gruber (2006), b Jin et al. (2006), c Millimann and Droxler (1996), d Feely et al. (2004), e Treguer et al. (1995)
changes and the response in terrestrial biosphere is not taken
into account. The simulation set consists of separate runs in
which SO wind stress is gradually varied between 20 % and
180 % relative to the original strength by steps of 20 %.
It has been found in several global ocean models (including ours) that deep mixing in the SO is driven to a large degree by momentum transfer from SH westerly winds to the
ocean surface in the latitudinal band spanned by the Drake
Passage (e.g., Toggweiler et al., 2006; Menviel et al., 2008;
Tschumi et al., 2008). Furthermore, the modeled characteristics of deep ocean water masses are affected on a global scale
by changes in SO deep mixing. Accordingly, the SO represents a major route of deep ocean ventilation in these ocean
models.
Yet, it remains hard to draw conclusions on the physical
mechanisms controlling deep mixing in the real SO. In recent studies, factors other than wind stress have been proposed to play a major role for the control of SO deep mixing.
These are for instance sea-ice dynamics and buoancy forcing (e.g., Fischer et al., 2010; Watson and Naveira-Garabato,
2006), small-scale eddy flows which are not represented explicitely in most global ocean models (e.g., Spence et al.,
2009; Böning et al., 2008) or other processes affecting the
vertical stratification in the SO (Sigman et al., 2004; Paillard
and Parennin, 2004). The relative importance of the different
mechanisms capable of driving glacial-interglacial variations
in SO deep mixing remains under debate. For this reason,
the amplitude of SO winds in our simulations is regarded as
a mere tuning knob for the circulation state in the model.
Clim. Past, 7, 771–800, 2011
2.3.1
Ocean circulation response
The two most prominent features in the circulation response
to SO wind variations are: (1) changes in the frequency and
depth of convective events in the SO and (2) changes in
the intensity and depth of the Atlantic meridional overturning circulation (AMOC). Stronger winds intensify deep convection and Ekman upwelling in the SO and strengthen the
AMOC (Fig. 5b and Tschumi et al., 2008). Weaker winds
induce the opposite response characterized by a more stably
stratified Southern Ocean and less intense AMOC.
In case of strongly reduced wind stress (20 % or 40 % of
the original amplitude), the AMOC breaks down completely
in the model, leaving the SO as the sole region of deep water formation. These latter circulation states are named “offstates” here, whereas the model states with active AMOC are
denoted as “on-states”.
As long as the AMOC remains active, i.e. in the on-states,
deep ocean ventilation scales quite homogeneously across
all ocean basins with the changes in SO wind amplitude
(Fig. 5a). Deep ventilation ages below 2000 m vary almost
linearly with the strength of SO wind stress. Globally averaged deep ventilation ages increase by 32 years per 10 %reduction in SO wind strength. The Atlantic deep water age
is slightly more sensitive (43 years/10 %-reduction) than in
the other basins, owing to the changes in AMOC strength.
In case of the off-states the link between deep ocean ventilation ages and SO wind stress deviates from the quasi-linear
relationship (Fig. 5a). In particular in the deep Atlantic,
the AMOC shutdown results in a large anomalous drop of
roughly 600 years in ventilation age. This model response
indicates the change in end-members of Atlantic deep water
subsequent to the collapse of AMOC.
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T. Tschumi et al.: Deep ocean ventilation, carbon isotopes, marine sedimentation and the deglacial CO2 rise
Fig. 5. (a) Ventilation age in different ocean basins as a function of the SO wind stress amplitude. Ventilation age is calculated
from simulated 14 R-signatures in the ocean below 2000 m depth.
14 R is treated here as an ideal age tracer in the sense that it is
transported only by ocean circulation and the atmospheric signature is held constant
(14 Ratm = 1.0).
Ventilation age is calculated
as τbasin = λ1 ln 14 Ratm /14 Rbasin where λ = 1/8267 yr−1 is the decay constant for 14 C and 14 Rbasin is the average 14 R-signature in
the respective basin. (b) Southern Ocean convection index and Atlantic Meridional Overturning Circulation (AMOC) versus SO wind
stress. The SO convection index is proportional to the number of
model grid cells in the SO participating in convective mixing.
To sum up, changing wind stress in our model proves itself as an effective method to gradually vary globally averaged deep water age in a range between 1580 years (180 %
SO wind stress) and 2100 years (20 % SO wind stress). Major structural shifts in the modeled global circulation pattern
occur only if SO wind stress is strongly reduced to 20 % or
40 % where AMOC breaks completely.
2.3.2
Response to sub-millennial processes
Figure 6 shows the time-dependent response for a set of informative carbon cycle parameters. The reaction of atmospheric CO2 and its isotopic composition to changes in deep
ocean ventilation is characterized by a major fast response
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unfolding during approximately the first 2 kyr and a minor
slow response over several 10 kyr (Fig. 6a, g and h). The fast
response arises from the reorganization of the ocean carbon
pumps, the long-term response is driven by the slow adaptation of sediment formation to the altered biogeochemical
state of the ocean. Atmospheric CO2 , δ 13 CCO2 and 114 CCO2
have reached a quasi-equilibrium after 2 kyr in the oceanonly model, whereas in the ocean-sediment model sedimentation processes affect atmospheric carbon on a time scale of
several 10 kyr.
Figure 7 shows the correlation between simulated deep
ocean ventilation age and the changes in atmospheric CO2 ,
δ 13 CCO2 and 114 CCO2 at the time slice 2 kyr after perturbing
SO wind stress forcing. The results unambiguously display
that better ventilated deep ocean watermasses drive an increase in atmospheric CO2 , whereas δ 13 CCO2 and 114 CCO2
decline. Model results of the ocean-sediment model at 2 kyr
are not substantially different from those of the ocean-only
model since the effect of changes in sedimentation has not
yet fully materialized at this early stage of the response.
Atmospheric CO2 and its isotopic signatures follow an
almost linear relationship with deep ocean ventilation age
for both without and with sediments as long as there is
no AMOC collapse. Linearized sensitivities for δ 13 CCO2
(0.08 ‰/100 years) and for 114 CCO2 (10 ‰/100 years)
are almost identical in both model settings (Fig. 7b
and c). Changes in atmospheric CO2 are slightly larger
(−8 ppm/100 years) at the 2 kyr time slice (Fig. 7a).
AMOC shutdowns significantly affect atmospheric CO2
and 114 CCO2 . The simulated changes for off-states clearly
deviate from the linear relationship as displayed in Fig. 7a
and c. The AMOC shutdown induces an anomalous drawdown in atmospheric CO2 of ∼10 ppm without sediments
and of ∼20 ppm with sediments. The anomalous rise in
114 CCO2 is of ∼10 ‰ in the ocean-only model and of
∼15 ‰ in the ocean-sediment model. Atmospheric δ 13 CCO2
on the other hand is rather unsensitive to the shutdown of
AMOC. There is no substantial deviation from linearity in
the relation between δ 13 CCO2 with the rate of deep ocean
ventilation when AMOC shuts down (Fig. 7b).
The biological pump strongly dominates the net CO2 response subsequent to wind stress changes in our model (see
also Tschumi et al., 2008). The effect of the solubility pump
is minor as simulated changes in (potential) deep ocean temperature (below 2000 m) are ∼0.5 ◦ C at most (Table 2). This
translates into a maximum effect on atmospheric CO2 of
roughly 5 ppm given the sensitivity of CO2 -solubility to temperature in seawater (e.g., Sigman and Boyle, 2000). On the
other hand, the imposed changes in ocean ventilation have
a strong impact on the marine biological carbon pumps. Enhanced (reduced) ocean ventilation implies increased (diminished) upwelling of nutrients, thus promoting (weakening)
biological export from the surface.
Table 2 indicates that organic matter export diagnosed at
2 kyr after the wind stress switch scales roughly linearly with
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T. Tschumi et al.: Deep ocean ventilation, carbon isotopes, marine sedimentation and the deglacial CO2 rise
Fig. 6. Ocean ventilation experiments: time-dependent carbon cycle response of the coupled ocean-sediment system to changes in deep
ocean ventilation induced by scaling the amplitude of windstress and deep convection in the Southern Ocean.
ventilation in case of the on-states (rising by 2.9 % per 100year reduction in global deep water age). The marine silica
cycle reacts substantially more sensitively to the ocean ventilation changes than the organic matter cycle: Opal export at
2 kyr increases by 8.6 % upon a 100 year-reduction in global
deep water age and decreases by 7.6 % per 100 year-increase.
The full time-dependent response in POC and opal export is
shown in Fig. 6c and d.
Clim. Past, 7, 771–800, 2011
Note that the efficiency of the soft-tissue pump P ∗, defined here as the ratio between regenerated and total phosphate in the deep ocean (Ito and Follows, 2005), behaves
conversely, i.e. P ∗ increases as export production decreases
parallel to slowed ventilation (Table 3). As long as NADW
formation is active, a linear relation is found between atmospheric CO2 and P ∗ (Fig. 8), which reflects a maximum impact of the biological pumps of 336 ppm at full efficiency.
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785
Table 2. Ocean ventilation experiments in the ocean-sediment model: simulated changes in the global mean deep ocean temperature,
biological production and in the effeciency of the soft-tissue pump at the time slice 2 kyr after the abrupt variation in ocean ventilation.
SO wind stress
amplitude
Atm.
pCO2
Mean deep
water agea
Mean deep
water temp.a
POM
export
CaCO3
export
Global mean
rain ratio
Opal
export
Efficiency of
soft-tissue
pumpb
%
ppm
years
◦C
GtC yr−1
GtC yr−1
–
Tmol Si yr−1
20 %
40 %
246.2
250.0
2097
2062
2.02
1.51
11.07
11.42
0.99
1.02
0.089
0.089
76.8
79.1
52.0
48.8
60 %
80 %
Control
120 %
140 %
160 %
18 0%
268.2
273.1
278.0
284.2
290.1
295.3
300.0
1963
1908
1840
1770
1706
1640
1578
1.49
1.54
1.61
1.62
1.69
1.77
1.90
12.14
12.54
12.76
12.99
13.24
13.46
13.62
1.08
1.11
1.13
1.14
1.16
1.17
1.18
0.089
0.089
0.088
0.088
0.088
0.087
0.087
86.4
92.7
95.5
98.6
104.2
110.7
116.4
47.7
46.7
45.4
43.5
41.7
40.0
38.4
a Global mean values below 2000 m depth.
b Defined as the fraction of regenerated PO versus total PO after Ito and Follows (2005), Eq. (4).
4
4
Fig. 7. Response in atmospheric (a) CO2 , (b) δ 13 CCO2 , and (c) 114 CCO2 to changes in ocean ventilation. Displayed values are diagnosed
for the time slice at 2 kyr after the perturbation in SO wind forcing. At this time carbon, cycling has reached a new equilibrium in the
ocean-only model (circles), whereas the ocean-sediment model (stars) is still under relaxation at 2 kyr due to the millennial-scale response
of ocean-sediment interactions.
This value is close to the 312 ppm found by Ito and Follows
(2005) on the basis of a simple theory combining the mass
balance for carbon and nutrients in the ocean-atmosphere
system and the linearization of carbonate chemistry. The relatively close correspondence between this theoretical value
and the sensitivity of atmospheric CO2 to P ∗ found in our
experiments indicates that the predominant part of the simulated CO2 response to changes in SO deep water formation
is due to changes in the biological pump efficiency.
The simulated relation between atmospheric CO2 and the
soft-tissue pump efficiency P ∗ deviates from linearity as
soon as NADW formation shuts down in response to weak
SO winds (Fig. 8). Thus, the anomalous CO2 reduction of
∼10 ppm due to the AMOC collapse cannot be explained
with changes in the biological carbon pumps. Further, the
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absence of a drop in simulated global mean deep water
temperature and in the global rain ratio (Table 2) seems
to preclude a dominant role of the solubility and carbonate
pumps. We conjecture that the simulated anomalous CO2
drop in response to collapsed AMOC results primarily from
an increase in the deep water DIC disequilibrium component
(Marinov et al., 2008).
The simulated interplay between ocean ventilation, marine
carbon cycling and atmospheric CO2 displayed in our wind
stress-experiments is in good qualitative agreement with the
results of Schmittner et al. (2007), who find a similar response in marine biogeochemistry to changes in the global
strength of wind stress. However, the effect of AMOC shutdowns on marine carbon cycling in our experiments cannot be readily compared with the corresponding results of
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T. Tschumi et al.: Deep ocean ventilation, carbon isotopes, marine sedimentation and the deglacial CO2 rise
2.3.3
Fig. 8. Ocean ventilation experiments: quasi-linear relation between simulated atmospheric CO2 and the efficiency of the softtissue pump P ∗ (Ito and Follows, 2005) for the ventilation states
with active NADW formation (circles). The linear relation is perturbed when NADW formation is shutdown in response to strongly
reduced windstress in the Southern Ocean (rectangles).
Schmittner et al. (2007), since AMOC is suppressed by freshwater forcing in the North Atlantic in their experiments rather
than by weak SO winds.
The reorganization of the marine biological cycle is further
reflected in the simulated change in the nutrient distribution
and δ 13 CDIC signatures. The total phosphate concentrations
decrease and δ 13 CDIC signatures increase in the sea surface
as ocean ventilation is slowed and euphotic zone nutrients are
utilized more efficiently (Table 3). After propagation of this
signal across the air-sea interface δ 13 CCO2 is shifted towards
less negative values in the atmosphere. The radiocarbon signature in atmospheric CO2 , 114 CCO2 , increases in response
to reduced ventilation due to slowed physical transport of radiocarbon to the deep ocean which ultimately reduces ocean
uptake of 14 C.
A specific consequence of the Southern Ocean ventilation
mechanism is that vertical δ 13 CDIC -gradients in the Southern Ocean should have steepened during glacials (Hodell
et al., 2003; Ninnemann and Charles, 2002). Our results reproduce this connection between deep mixing in the Southern Ocean and the vertical gradient in δ 13 CDIC . Reduced
deep mixing evidenced by a larger 114 CDIC difference between the surface and deep Southern Ocean is accompanied by steeper vertical gradients in total phosphate and
δ 13 CDIC (Table 3). According to our results, upon a reduction of 100 years in the difference between surface and
deep water age (mean value below 2000 m) in the Southern Ocean, the δ 13 CDIC -gradient (difference between surface
and mean signature below 2000 m in the Southern Ocean) increases by 0.16 ‰ and the vertical phosphate-gradient grows
by 0.13 mol m−3 (Table 3).
Clim. Past, 7, 771–800, 2011
The long-term effect of ocean-sediment
interactions
The long-term millennial-scale trends in atmospheric CO2 ,
δ 13 CCO2 and 114 CCO2 arise from the slow response in sediment diagenesis to the changing biogeochemical state of the
ocean. The sedimentary processes tend to restore in the long
run burial rates of CaCO3 , POC and opal towards initial unperturbed values in order to re-establish the balance with the
fixed weathering inputs.
The simulated variations in atmospheric CO2 are amplified by 30 %–60 % within 50 kyr through the net effect of
the sedimentary feedbacks subsequent to ventilation changes
(Table 4). The overall simulated sedimentary CO2 response
is dominated by the changes in CaCO3 burial. In fact, the
shifts in organic matter burial counteract the simulated CO2
variations as long as AMOC is maintained (on-states). For
instance, accelerated ocean ventilation leads to a net rise
in atmospheric CO2 , whereas elevated organic matter burial
tends to reduce CO2 by drawing carbon out of the oceanatmosphere system. In the on-states, the sedimentary CO2
response amplification is of 30 %–40 % (Table 4). In case of
collapsing AMOC on the other hand, CO2 changes are reinforced by the sedimentary feedbacks by as much as 60 %,
since both carbonate compensation from reduced CaCO3 export as well as excess POC burial driven by the reduction in
deep ocean oxygen availability reinforce the CO2 drawdown
from slowed ventilation.
In contrast to the previous experiment in which the wholeocean nutrient content is enhanced, the concomitant variations in deep ocean oxygen availability subsequent to ventilation changes weaken the importance of the nutrient-feedback
by decoupling the link between shifts in POC export and
in POC sediment burial. Enhanced (decreased) ventilation
promotes (reduces) remineralization within the sediments
through increased (decreased) oxygen concentrations in pore
waters. A specific consequence is that the changes in POC
export arising from altered deep ocean ventilation are maintained to a substantial degree in the long term (Fig. 6c). The
POC burial flux, however, relaxes again to initial conditions
to re-establish the balance with weathering input. In conclusion, the variations in oxygen availability in the ocean
driven by the ventilation changes significantly decouple the
link between POC export and POC burial. This decoupling
complicates the reconstruction of past POC export on the basis of sedimentary records.
The sedimentary response in atmospheric δ 13 CCO2 is dominated by the temporal evolution of POC burial. For instance,
in the model states with accelerated ocean ventilation, excess
POC burial draws organic matter with low δ 13 (∼ −20 ‰)
out of the ocean-atmosphere system. As a consequence, the
sediment response is to shift δ 13 CCO2 towards less negative
values (Fig. 6g, red and yellow lines). Excess burial of POC
and CaCO3 both remove carbon with low 114 C from the
ocean-atmosphere system in response to a better ventilated
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Table 3. Ocean ventilation experiments in the ocean-sediment model: simulated changes in ocean biogeochemistry of different ocean regions
at 2 kyr.
SO wind stress
amplitude
North Atlantic below 2000 m
(0◦ –71◦ N)
reg
North Pacific below 2000 m
(0◦ –63◦ N)
114 C
[‰]
δ 13 C
[‰]
PO4
[µmol l−1 ]
PO4
[µmol l−1 ]
O2
[µmol l−1 ]
114 C
[‰]
δ 13 C
[‰]
PO4
[µmol l−1 ]
20 %
40 %
−260
−258
0.171
0.179
2.54
2.55
1.31
1.23
118.2
134.7
−273
−265
0.092
0.011
60 %
80 %
Control
120 %
140 %
160 %
180 %
−195
−188
−180
−173
−168
−163
−156
0.632
0.671
0.711
0.734
0.745
0.754
0.774
1.98
1.94
1.88
1.84
1.82
1.78
1.74
0.88
0.85
0.81
0.77
0.73
0.70
0.66
191.0
194.8
198.9
203.3
205.9
208.5
211.9
−279
−279
−278
−276
−274
−273
−272
0.011
0.026
0.049
0.077
0.102
0.123
0.141
SO wind stress
amplitude
Southern Ocean below 2000 m
(South of 51◦ S)
reg
O2
[µmol l−1 ]
2.61
2.60
1.37
1.26
106.4
129.1
2.74
2.73
2.70
2.68
2.66
2.64
2.62
1.33
1.31
1.28
1.22
1.17
1.13
1.10
117.5
121.4
127.1
136.7
144.6
151.4
156.9
Southern Ocean surface
(South of 51◦ S)
114 C
[‰]
δ 13 C
[‰]
PO4
[µmol l−1 ]
PO4
[µmol l−1 ]
O2
[µmol l−1 ]
114 C
[‰]
δ 13 C
[‰]
δ 13 CS−D
[‰]
20 %
40 %
−196
−198
0.385
0.377
2.27
2.30
1.06
0.98
160.8
178.9
−61
−73
2.350
2.198
60 %
80 %
Control
120 %
140 %
160 %
180 %
−208
−208
−207
−205
−204
−202
−199
0.308
0.330
0.358
0.392
0.425
0.455
0.490
2.40
2.39
2.36
2.34
2.32
2.30
2.27
1.01
0.98
0.94
0.88
0.82
0.77
0.72
175.3
180.1
186.5
197.2
206.1
214.5
222.3
−99
−106
−112
−119
−125
−130
−135
2.049
1.944
1.871
1.799
1.725
1.658
1.603
deep ocean. As a result, atmospheric 114 CCO2 is driven towards slightly less negative values (Fig. 6h, red and yellow
lines).
Figure 6b (red and yellow lines) displays that an accelerated ocean ventilation further induces an initial deepening in
the CaCO3 saturation horizon which is offset after 1–2 kyr.
The first part of the response results from the removal of carbon from the ocean interior as a result of rejuvenated deep
water masses and excess POC sedimentation. The subsequent shoaling is due to the slowly responding effect of carbonate compensation to excess CaCO3 burial.
The model experiments with collapsing AMOC exhibit the
strongest reduction in nutrient supply to the surface and in
globally integrated POC export from the surface (blue line in
Fig. 6c). Global POC export remains reduced significantly
in the long run. However, subsequent to the strong initial
decrease, POC export in the North Atlantic eventually increases after about 200 years driven by diffusion of nutrients along isopycnals to the North Atlantic ocean surface.
This response in POC export production is qualitatively similar to the results of Chikamoto et al. (2008) where AMOC
has been shut down through a freshwater discharge in the
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reg
PO4
[µmol l−1 ]
PO4
[µmol l−1 ]
O2
[µmol l−1 ]
1.964
1.821
0.57
0.69
297.3
296.9
1.741
1.614
1.513
1.407
1.301
1.203
1.113
1.04
1.11
1.18
1.24
1.31
1.37
1.41
296.7
296.5
295.8
295.4
295.0
294.6
294.5
North Atlantic. In our model, we additionally find a longterm decreasing trend in global POC export after 5 kyr which
is due to whole-ocean nutrient loss by excess POC burial.
This nutrient-feedback has not been taken into account by
Chikamoto et al. (2008).
Similarly to POC export, organic matter burial strongly declines during the first 200 years (blue line in Fig. 6e). Thereafter, the burial rate sharply recovers as a result of increasing POC export in the North Atlantic and a strongly reduced
deep ocean oxygen supply. After around 2500 years, there
is excess POC burial in spite of the reduction in global POC
export production. Opal export and opal sediment burial exhibit a similar response as POC export and burial do (Fig. 6d
and f): The AMOC shutdown induces major shifts also in
the spatial pattern of opal export and sedimentation, with the
North Atlantic becoming a region of deep nutrient upwelling
similar to the North Pacific.
The reorganization in the marine biological carbon cycle
in case of AMOC shutdowns is further reflected in the timedependent model response for the depth of the CaCO3 saturation horizon. The AMOC collapse induces an initial shoaling of the CaCO3 saturation horizon as DIC accumulates in
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T. Tschumi et al.: Deep ocean ventilation, carbon isotopes, marine sedimentation and the deglacial CO2 rise
Table 4. Long-term response in atmospheric CO2 and amplification by sedimentary feedbacks.
SO wind stress
amplitude
Global Deep
Water Age
[years]
Atmospheric CO2
without sediments
after 50 kyr
[ppm]
Atmospheric CO2
with sediments
after 50 kyr
[ppm]
Sedimentary
amplification
of the CO2 response
[%]
20 %
40 %
2097
2062
251
252
234
236
63 %
62 %
60 %
80 %
Control
120 %
140 %
160 %
180 %
1963
1908
1840
1770
1706
1640
1578
269
273
278
285
290
295
299
265
271
278
287
294
302
308
44 %
40 %
–
29 %
33 %
30 %
30 %
the deep from the slowed ventilation and reduced POC burial
(Fig. 6b). Consistently, the shoaling is more pronounced in
the Atlantic than in the Pacific basin. After 1–2 kyr the saturation horizon starts to deepen again in both basins as deep
[CO2−
3 ] rises from the combined effects of reduced CaCO3
burial and increasing POC burial. The final steady state shift
in the CaCO3 saturation horizon is a deepening of 100–200 m
in the Atlantic and of 500–600 m in the Pacific.
3
Discussion
Our modelling approach is associated with a number of conceptual and methodological limitations worth mentioning.
First, the model scenarios are idealized and not intended to
directly reflect the temporal evolution of paleo-proxy signals
as recorded in climatic archives. Second, changes in deep
ocean ventilation are generated by varying the magnitude of
the wind stress in the Southern Ocean by more than a factor
of two. This is a convenient technical tuning mechanism to
adjust the deep ocean ventilation age in the model. However,
we stress the fact that these variations in SO wind strength
exceed the range of realistic changes. The atmospheric CO2
response to realistic variations has been found to be rather
modest (Tschumi, 2009; Menviel et al., 2008).
The model simulations have been tailored to investigate
the specific response in atmospheric CO2 and other paleoproxies to uniform variations in deep ocean ventilation age.
Several possibly important processes to explain the full
glacial-interglacial CO2 variability have not been taken into
account. We have neglected, for instance, the well-recorded
variations of AMOC (McManus et al., 2004), changes in the
terrestrial biosphere, in sea surface temperatures or in sea ice
cover. Further, the Bern3D+C is a model of intermediate
complexity with coarse resolution and does not resolve eddies and other processes operating on small scales. Marine
Clim. Past, 7, 771–800, 2011
biological production is represented as simple prognostic formulations considering the competition between opal and calcite producers and changes in the uptake ratio of silica to
phosphate.
When interpreting the model results in the context of carbon cycle changes during the last deglaciation it should be
kept in mind that the starting point of all sensitivity experiments is a stationary state corresponding to the preindustrial
climate and carbon cycle. This modelling approach has been
chosen, as the knowledge about glacial climate is subject to
considerable uncertainty. The focus of this study is further
laid upon a more qualitative interpretation of the interplay
between different carbon cycle processes which depends less
critically on the specific choice of initial conditions than the
quantitative interpretation of a single paleoclimatic process
or signal.
Nevertheless, we have have performed an additional sensitivity experiment in order to quantify the uncertainty associated with the initial state of ocean ventilation: Starting from
a stationary state with preindustrial carbon cycle boundary
conditions but with low ocean ventilation as resulting from
60 %-SO winds, SO wind stress has been increased to 140 %.
Comparing the response in atmospheric CO2 , δ 13 CCO2 and
114 CCO2 with the corresponding 180 %-SO winds experiment starting from standard winds (red lines in Fig. 6) has
shown that the sensitivity to the initial ventilation state of
the ocean is negligible as long as there is no structural shift
in the large-scale circulation pattern such as a breakdown of
AMOC.
In spite of the inherent limitations mentioned above, the
results of our study allow for a number of robust conclusions.
We argue in the following that the processes represented in
our model played a dominant role for reorganizing the global
carbon cycle during the early part of the last deglaciation.
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T. Tschumi et al.: Deep ocean ventilation, carbon isotopes, marine sedimentation and the deglacial CO2 rise
3.1
Southern Ocean ventilation hypothesis
The “Southern Ocean ventilation hypothesis” (e.g., Francois
et al., 1997) is the starting point for discussion of our modelling results. According to this hypothesis, low glacial CO2
values are related to a poorly ventilated deep ocean and a stably stratified Southern Ocean. Such a configuration would
lead to an accumulation of carbon in the deep ocean during glacial periods relative to interglacials. The rise in atmospheric CO2 over the glacial transition is driven by the breakup of Southern Ocean stratification and the resumption of
modern-scale upwelling, bringing water rich in carbon, silica, and other nutrients and depleted in 13 C and 14 C to the
surface. The characteristics of these upwelled water masses
are further communicated within the intermediate and surface ocean by entrainment into the Subantarctic Mode Water
and Antarctic Intermediate Water and to the atmosphere by
air-sea gas-exchange. During recent years, a large number
of observational evidence has been published supporting the
view that changes in SO deepwater formation and in deep
ocean ventilation have played an important role for glacialinterglacial CO2 variations (e.g., Fischer et al., 2010; Spero
and Lea, 2002; Adkins et al., 2002; Anderson et al., 2009;
Muscheler et al., 2004; Brovkin et al., 2007; Hodell et al.,
2003; Marchitto et al., 2007; Francois et al., 1997).
The response of our ocean-sediment carbon cycle model
to strengthened Southern Hemisphere westerly winds reproduces some of the key mechanisms attributed to the “Southern Ocean ventilation hypothesis”. Invigorated wind stress
promotes vertical exchange of water masses in the SO,
thereby significantly affecting the physical and biogeochemical properties of deep ocean water on a global scale. Our
results highlight in particular a rather close correlation between changes in ventilation age, the amount of biologically
sequestered carbon and the δ 13 C-signature in sea water as
shown in Fig. 9.
In the following sub-sections we will discuss paleo-data
from sediment analyses and from ice cores which indicate that large changes in Southern Ocean ventilation indeed occurred during the early part of the deglaciation,
roughly in the interval between 18 ka BP and the onset of
the Antarctic Cold reversal and the Bølling/Allerød Northern Hemisphere warming at about 14.6 ka BP (Severinghaus
and Brook, 1999). We will then compare the data with our
model results to constrain the magnitude of the change in atmospheric CO2 over the last termination which is driven by
a SO ventilation mechanism.
3.1.1
Changes in Southern Ocean opal sedimentation
Anderson et al. (2009) report reconstructed opal fluxes in
Southern Ocean sediments which show a rapid increase starting roughly together with the rise in atmospheric CO2 around
17 ka BP, followed by a deglacial maximum and a long-term
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789
decrease into the Holocene (Fig. 9). Opal flux rates are
typically several-fold higher during the transition compared
to glacial and late-Holocene values. Anderson et al. (2009)
interpret these opal flux changes as variations in Southern Ocean upwelling, concluding that there was a rapid increase in the upwelling of silica and carbon rich deep water to the Southern Ocean surface at the onset of deglaciation (∼18 ka BP). They propose a direct link between the increased ventilation of deep water to the deglacial rise in atmospheric CO2 .
The step-wise strengthening of SO wind stress as applied
in our ocean ventilation experiments drives an immediate rise
in deep upwelling in the Southern Ocean in response to enhanced Ekman pumping and more vigorous deep convection
in the model. These simulations may thus serve as model
scenarios for an abrupt resumption of upwelling in the SO.
Figure 9 shows that our model results for strengthened SO
wind stress are in good qualitative agreement with the sediment data of Anderson et al. (2009) in relative magnitude
and timing: Increasing SO wind stress by 80 % drives a rapid
30 %-rise in opal export as more silica is supplied to the surface. Additionally, opal sedimentation also increases and
reaches a peak value with a delay of several millennia relative to the reinvigoration of deep upwelling in the Southern
Ocean. The peak sedimentation flux is two to three times
higher than the initial flux. Both opal export and sedimentation slowly decrease towards initial rates with a typical
time scale of 20 kyr. Note however that simulated opal burial
fluxes are about 10 times smaller than the reconstructed opal
flux (Fig. 9). This is likely due to the spatial averaging over
the model grid cell and the relatively low simulated global
burial rate of opal as compared to available estimates (Table 1).
Anderson et al. (2009) suggest, based on the positive correlation between 231 Pa/230 Th and opal flux in their records,
that the deglacial maximum in opal flux reflects past changes
in opal production and Si supply rather than variable opal
preservation. In our model however, a relatively modest increase in global silica export of 30 % enhances the net flux to
the sediments in a highly non-linear way as pore water concentrations of silicic acid and thus the preservation of opal
in the sediments greatly increases under the enhanced rain of
silica particles (Archer et al., 2000a; Ridgwell et al., 2002).
The majority of the simulated change in opal sedimentation
flux thus reflects increased preservation rather than increased
opal production as suggested by Anderson et al. (2009).
Why does the reconstructed opal sedimentation flux relax
towards glacial conditions over the course of the Holocene,
whereas atmospheric CO2 and, presumably, Southern Ocean
upwelling remains at high interglacial values? From a continued large Southern Ocean upwelling, one might expect a
continued silica supply to the surface ocean, and a sustained
high export and sedimentation rate. In our simulation, increased opal burial drives a loss of whole-ocean silica. As a
result, the simulated export of opal and its sedimentation rate
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T. Tschumi et al.: Deep ocean ventilation, carbon isotopes, marine sedimentation and the deglacial CO2 rise
Fig. 9. Model response for δ 13 CDIC , ventilation age and 114 CDIC to stimulated deep mixing in the Southern Ocean and to more vigorous
and deeper AMOC as a result of a 80 %-increase in SO wind stress amplitude. Shown are the simulated differences in Atlantic and Pacifc
zonal mean sections at 15 ka after the the step-wise perturbation of the model state.
relax towards initial conditions to balance riverine input with
the burial flux of silica.
Is the evolution of the opal sedimentation flux directly correlated with opal export and the resumption of upwelling or
are there any system time lags? The simulated peak in opal
sedimentation is delayed by several millennia compared to
the onset of the ventilation in our model. This suggests that
Clim. Past, 7, 771–800, 2011
the temporal evolution of the ventilation history may be difficult to disentangle exactly from the temporal evolution of the
opal sedimentation flux. Nevertheless, the opal records appear consistent with a change in deep ocean ventilation that
occurred during the early deglacial period and that may or
may not have been abrupt.
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T. Tschumi et al.: Deep ocean ventilation, carbon isotopes, marine sedimentation and the deglacial CO2 rise
3.1.2
Changes in the stable carbon isotope ratio δ 13 C
Figure 10 displays the currently available δ 13 CCO2 records
from Antarctic ice cores covering the period of the last
deglaciation (Smith et al., 1999; Lourantou et al., 2010;
Elsig, 2009) together with reconstructed atmospheric CO2
(Monnin et al., 2001) and δ 18 OH2 O (Stenni et al., 2001),
which is a proxy for local temperature on the ice sheet. The
δ 13 CCO2 records consistently show an early deglacial decrease of 0.3–0.4 ‰ to reach a minimum at around 15 ka BP
and a millennial-scale long-term increase by the same magnitude to attain approximately constant values at 7 ka BP (Elsig
et al., 2009). However, the records disagree on the details of
the δ 13 CCO2 evolution during the Bølling/Allerød-Younger
Drias climate swings.
Sediment cores from different ocean basins (Indo-Pacific,
sub-Antarctic, and South Atlantic) reveal that the occurrence of carbon isotope minima at the beginning of glacial
terminations is a widespread feature in planktic foraminifera
δ 13 C records (Spero and Lea, 2002). The decrease in the
δ 13 C records occurred simultaneously with the initiation of
Southern Ocean warming. The data suggests a decrease in
surface water δ 13 C in the early deglacial period (∼20 ka BP
to 16 ka BP), an absolute minimum at the same time as the
ice core record of Smith et al. (1999) which is followed by
an increase over the rest of the deglacial period into the early
Holocene. Taken together, both sediment and ice core data
suggest decreasing atmospheric δ 13 C during the first stage
of the deglacial period followed by a long-term increase of
similar magnitude.
How can this millennial-scale δ 13 C evolution be explained? Our results lend support for the interpretation of
Spero and Lea (2002) and Köhler et al. (2005) that the early
deglacial δ 13 C decrease reflected renewed Circumpolar Deep
Water upwelling with an enhanced transport of low δ 13 C. A
decrease in atmospheric δ 13 CCO2 of 0.3 ‰ would require a
rejuventation of deep ocean waters by 300 to 400 years according to the δ 13 CCO2 sensitivity found in our ventilation
experiments (Fig. 7b). Subsequently, this negative signal is
reduced by roughly 50 % over the following 15 kyr due to excess sedimentation of organic material (Fig. 6g). This would
leave an increase of 0.1–0.2 ‰ to be explained by other processes such as carbon uptake by the terrestrial biosphere, surface ocean warming or sea ice retreat.
3.1.3
791
Changes in CaCO3 saturation horizon depth
Analyses of the sedimentary bulk composition indicate a period of enhanced preservation of CaCO3 deep sea sediments
between ∼15–7 ka BP (Ridgwell et al., 2003; Broecker
et al., 2001). Broecker et al. (2001) attribute this sedimentary signal to an early deglacial restoration of the terrestrial biosphere. The resulting removal of carbon from the
atmospheric and oceanic reservoirs is made responsible for
a deepening of the CaCO3 saturation horizon between about
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Fig. 10. Reconstructed local opal flux (blue line) at 61.9◦ S/170◦ W
from Anderson et al. (2009) and simulated opal burial rate in the
corresponding model cell (red line) as simulated in the experiment
with an abrupt 80 %-increase in SO wind stress at 18 ka BP.
15–10 ka BP. The deepening is followed by a gradual shoaling until the latter part of the Holocene which is attributed
to carbonate compensation. Ridgwell et al. (2003) invoke regrowth of coral reefs as additional factor driving the deglacial
signal of the Pacific CaCO3 saturation horizon to simultaneously account for the evolution of atmospheric CO2 .
In response to stimulated upwelling in the Southern Ocean
rejuvenating deep waters by 300–400 years, our model predicts a similar temporal evolution for the Pacific CaCO3 saturation horizon. We simulate a deepening saturation horizon
by 300–400 m during the first ∼2 kyr and a slow shaoling
thereafter to reach initial depths after around 28 kyr (Fig. 6b).
The simulated initial deepening is a result of the flushing of
the old carbon-rich deep water mass from the Pacifc basin
and of excess organic matter sedimentation (Fig. 4b). This
is consistent with deglacial CaCO3 and organic matter flux
reconstructions from the North Pacific (e.g., Jaccard et al.,
2005; Galbraith et al., 2007) as well as with CO32− reconstructions based on benthic Zn/Ca measured in the deep
equatorial Pacific (Marchitto et al., 2005). However, the timing of our model results does not match the observations
very precisely, maybe due to the unrealistic abrupt change in
model boundary conditions. Nevertheless, the agreement is
sufficient in order to indicate that increasing deep ocean ventilation and early deglacial excess sedimentation of organic
matter may have significantly contributed to the recorded
carbonate ion signal in the Pacific ocean in addition to the
changes in terrestrial biosphere and growth of coral reefs
as proposed by Broecker et al. (2001) and Ridgwell et al.
(2003).
3.1.4
Changes in radiocarbon signatures
Is an increase in deep ocean ventilation consistent with marine radiocarbon data? To generate the full amplitude of
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792
T. Tschumi et al.: Deep ocean ventilation, carbon isotopes, marine sedimentation and the deglacial CO2 rise
the early deglacial 35 ppm increase in atmospheric CO2 , a
decrease in the deep ocean ventilation age of about 300–
400 years is required according to the model-derived linear relationship between CO2 and ventilation age (Fig. 7a).
Corresponding simulated changes in deep ocean 114 CDIC
are an increase of 40-45 in the deep North Atlantic of
10–15 ‰ in the deep North Pacific and in the deep Southern Ocean (Table 4). Atmospheric 114 C would drop by
30–40 ‰ (Fig. 7c). The simulated atmosphere-deep ocean
114 C-difference would thus be reduced by 70–80 ‰ in the
deep Atlantic and by 40–50 ‰ in the deep North Pacific and
the deep Southern Ocean.
A decrease in deep ocean ventilation age of 300–400 years
during the early deglacial period seems broadly consistent
with available proxy records. However, the locally large
changes between 18 ka BP and 14.5 ka BP in the benthicplanktonic age-difference of ∼1000 years in the Atlantic sector of the SO (Skinner et al., 2010) or of ∼700 years in the
deep North Pacific (Galbraith et al., 2007) are not reproduced
in our model. Overall, our results are in line with the modelbased interpretation of the glacial-interglacial changes in past
atmospheric 114 C by Köhler et al. (2006b) where a ∼40 ‰contribution is attributed to changes in ocean circulation,
compared to the 30–40 ‰ from our results.
A difficulty in the interpretation of marine 114 C data in
terms of ventilation age is that the 114 C signature at any
given time does not just reflect the ventilation age, but also
the time history of the atmospheric 114 C. The life time of
radiocarbon with respect to radioactive decay is 8267 years,
implying a considerable memory to earlier 114 C variations.
Atmospheric 114 C decreased by about 400 ‰ over the last
transition (Hughen et al., 2004; Beck et al., 2001). A riddle
presents the as yet unexplained drop of about 190 ‰ between
17.5 to 14.5 ka BP (Broecker and Barker, 2007). This drop is
almost four times larger than the scenario-based decrease of
50 ‰ and apparently also finds no correspondence in deep
ocean radiocarbon change (Broecker and Barker, 2007). We
are cautious to draw conclusions at this stage and believe that
further evaluations are required using transient simulations of
the coevolution of atmospheric CO2 , 114 C, and 14 C productivity.
4
Conclusions
Can we estimate the contribution of the Southern Ocean ventilation mechanism to the last deglacial CO2 rise on the basis of our model results? CO2 rose by about 35 ppm between the LGM and 15 ka BP (Fig. 10). If this CO2 increase is entirely attributed to rejuvenating deep ocean waters, it would translate into a global reduction in deep water
age of roughly 350 years, according to the model-derived linear relationship between CO2 and ventilation age (Fig. 7a).
Due to carbonate compensation in response to this ventilation change, CO2 would increase by another ∼10 ppm over
Clim. Past, 7, 771–800, 2011
Fig. 11. Paleo-records from Antarctic ice cores: δ 13 C-data is from
Elsig et al. (2009) (green triangles), Smith et al. (1999) (red triangles) and Lourantou et al. (2010) (blue triangles). The blue arrow indicates the recorded drop in δ 13 CCO2 (∼ −0.4 ‰) starting
around 18 ka BP. CO2 -data is from Monnin et al. (2001) and δ 18 Odata from Stenni et al. (2001).
the following 15 kyr. The early deglacial CO2 increase might
also reflect positive contributions from surface water temperature changes (Tschumi et al., 2008) and an estimated negative contribution from terrestrial uptake due to CO2 fertilization on ice-free land (Joos et al., 2004). The two effects
might have roughly cancelled. Taken together, our model results suggest that the Southern Ocean ventilation mechanism
has contributed about half to the reconstructed total glacialinterglacial CO2 increase of 90–100 ppm.
Our interpretation is further supported by a number of
prominent paleo-proxy signals that are adequately reproduced in our simulations. As a response to an abrupt rejuvenation of deep ocean watermasses by 300–400 years stimulated by more vigorous deep convection in the Southern
Ocean, we simulate a decline in δ 13 CCO2 by ∼0.3 ‰ during
the first 2 kyr followed by a gradual rise as observed in the
paleo-records. Taken at face value, the ice core data suggest
a slope of 12 ppm per 0.1 ‰ change in δ 13 CCO2 (35 ppm per
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T. Tschumi et al.: Deep ocean ventilation, carbon isotopes, marine sedimentation and the deglacial CO2 rise
0.3 ‰) during the early deglacial period, whereas the modelderived slope amounts to 13 ppm per 0.1 ‰. The model response also agrees well with increased opal sedimentation
rates in the Southern Ocean during the early deglaciation
(Anderson et al., 2009) and is further consistent with the reconstructed evolution of the Pacific CaCO3 saturation horizon (Broecker et al., 2001). We finally predict a drop of 30–
40 ‰ in atmospheric 114 CCO2 as another result of the invigorated deep ocean ventilation.
Although the simulated model response fits rather well
with these specific paleo-proxy records, a perfect match
should not be expected. Modeled carbon cycle sensitivities might be subject to considerable uncertainty, in particular due to unsettled questions related to the diffusivity of
the oceans (Archer et al., 2000b). Further, it turned out difficult to reproduce the exact temporal evolution of δ 13 C and
114 C, as recorded in sediment cores in the model. These
signals might sensitively depend upon regional and temporal
details of the deglacial reorganization of global ocean circulation and may be further affected by the variations of AMOC
(McManus et al., 2004) not considered here. Nonetheless,
based on the multi-proxy comparison between our simulated
response and the available paleo-data, we conclude that accelerating deep ocean ventilation linked to physical processes
in the SO has been the dominant driver of early deglacial
changes in global carbon cycling. Early deglacial ocean ventilation changes have continued to significantly contribute
to variations in atmospheric CO2 until the latter part of the
Holocene, since it takes millennia for marine sedimentation
to re-establish equilibrium with weathering input.
The glacial-interglacial CO2 variations appear to be driven
by a complex spatio-temporal pattern of processes involving relaxation time-scales of years to millennia. Our analysis of the time-dependent response in the coupled oceansediment system highlights that the benefit of conventional
“time slice” approaches to study these processes might be
rather limited. Our model results suggest that proxy signals arising from deglacial ventilation changes are sustained
on a multi-millennial time-scale until the balance between
carbon, nutrient and isotope input through rivers and winds
and the loss through sedimentation is re-established. Thus,
differences in these proxies between the LGM and the late
Holocene time slice may partly reflect imbalances in the
weathering-sedimentation cycles. This may also be relevant
for proxies not included here such as silica or nitrogen isotopes.
We have identified a feedback between the marine nutrient
inventory and sedimentation of organic matter. This mechanism emerges from the fact that marine biological productivity in the long run is limited by the riverine input of nutrients into the ocean. Our results indicate that this nutrientsediment feedback might be important for glacial-interglacial
changes in global carbon cycling. For instance, invigorated ocean circulation promotes export production and excess sediment burial of organic matter. The implications for
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793
the combined ocean-atmosphere carbon cycle are similar to
those of a postulated early deglacial restoration of the terrestrial biosphere (Broecker et al., 2001).
In this study, we have analyzed the time-dependent response in atmospheric CO2 and its isotopic signatures
only to the specific mechanisms of deep ocean ventilation
changes and biological pump reorganizations. To improve
the process-related understanding of the CO2 , δ 13 CCO2 and
114 CCO2 records over the last deglaciation, further mechanisms such as sea surface temperature changes, coral reef
growth, changes in the carbon stock on land and the spatiotemporal details of the various regional oceanic reorganizations should be considered in coupled 3-D global carbon cycle models. Simultaneous model-data comparison on
the basis of multiple proxy records wisely chosen seems to
be a promising approach towards a better understanding of
the dynamical interplay between the multiple factors driving
glacial-interglacial CO2 fluctuations. A particular challenge
is to identify and simulate the physical mechanisms responsible for the changes in Southern Ocean ventilation.
Appendix A
Sediment diagenesis model
The sediment diagenesis component used in this study is
built on previous work by Heinze et al. (1999) and Gehlen
et al. (2006) where the basic modelling concept is based on
Archer et al. (1993). The spatial model domain is restricted
to the diagenetical zone of the sediments which is assumed
here to be the top 10 cm of the surface sediments. Any solid
material leaving this domain disappears into the subjacent
diagenetically consolidated zone. The interfaces of the 10
model layers are defined at depths of 0, 0.3, 0.6, 1.1, 1.6, 2.1,
3.1, 4.1, 5.1, 7.55, and 10.0 cm. The fraction of pore water
volume (porosity) decreases from 0.95 to 0.75, following observations from Ullman and Aller (1982). Model geometry
remains fixed during integration.
The vertical movement of solid material in the sediment column is calculated on the basis of the equilibrium:
burial = rain redissolution. When the particulate rain onto
the sediment is larger than the redissolution of solid material within the diagenetical zone, the entire column is shifted
downwards and sediment burial occurs. Conversely, whenever redissolution is larger than rain, gaps of solid material
are produced within the model geometry. These are filled
with clay from below. Lateral movements in the sediments
are not considered.
Modeled tracers are the four solid components (CaCO3 ,
opal, POM and clay) and the six pore water substances (DIC,
total alkalinity, phosphate, nitrate, oxygen and silicic acid).
CaCO3 is only considered in the form of calcite. The less
stable aragonite form is assumed to be only of secondary importance for the global sediments (Berner and Honjo, 1981).
Clim. Past, 7, 771–800, 2011
794
T. Tschumi et al.: Deep ocean ventilation, carbon isotopes, marine sedimentation and the deglacial CO2 rise
Table A1. Parameters for the sediment component of the Bern3D+C model along with standard values and references.
Symbol
Parameter
Value
Reference
ρ
MPOM
MCaCO3
Mopal
Mclay
Number of time steps per year
Density of solid sediment components
Molar weight of POM
Molar weight of CaCO3
Molar weight of opal
Molar weight of clay
Terrestrial clay flux
48
2.6 g cm−3
32.74 g mol−1
100.0 g mol−1
67.2 g mol−1
430.51 g mol−1
1 g m−2 yr−1
this study
Heinze et al. (1999)
Heinze et al. (1999)
Heinze et al. (1999)
Heinze et al. (1999)
Heinze et al. (1999)
Heinze et al. (1999)
rCaCO3
ropal
roxy
rdenit
[Si]sat
Rate constant for CaCO3 redissolution
Rate constant for opal redissolution
Rate constant for oxydation
Rate constant for denitrification
Saturation concentration
for pore water silicic acid
1000 l mol−1 yr−1
20 l mol−1 yr−1
50 l mol−1 yr−1
50 l mol−1 yr−1
800 µmol l−1
Heinze et al. (1999)
this study
Heinze et al. (1999)
Gehlen et al. (2006)
Heinze et al. (2003)
D
DB
[O2 ]crit
Pore water diffusion coefficient
Bioturbation diffusion coefficient
Critical oxygen concentration
for denitrification
8. × 10−6 cm2 s−1
15. × 10−3 cm2 yr−1
1 µmol l−1
Heinze et al. (1999)
Heinze et al. (1999)
Gehlen et al. (2006)
RPO4
RO2
R−
NO3
RC
Rdenit
Redfield coefficient for phosphate
Redfield coefficient for oxygen
Redfield coefficient for nitrate
Redfield coefficient for carbon
Redfield coefficient for denitrification
1
170
16
117
97.6
Anderson and Sarmiento (1994)
Anderson and Sarmiento (1994)
Anderson and Sarmiento (1994)
Anderson and Sarmiento (1994)
Gehlen et al. (2006)
ρ
Mopal
All the solid components are taken to have equal densities
(ρ = 2.6 g cm−3 ).
Considered processes occurring during sediment diagenesis are the dissolution of opal and CaCO3 as well as the oxydation and denitrification of POM. The mixing of sediments
by burrowing animals such as clams and worms (bioturbation) is parametrized as a diffusion of solid components with
closed boundaries at the sediment-water interface and at the
bottom of the bioturbated zone.
The mathematical formulation of the sediment model consists of two sets of dynamical equations, one for the four
solid components and one for the six pore water substances,
along with adequate boundary conditions. The two sets of
equations are coupled to each other through the four reaction
rates of the sediment-pore water processes (redissolution /
remineralization). All of these reactions follow first order
kinetics with spatially uniform reaction rate constants. The
model parameters of the sediment component are listed in
Table A1 along with their standard values and references.
Ropal = ropal · copal · 8 · (1 − 8) ·
A1
RCaCO3 = rCaCO3 · cCaCO3 · 8 · (1 − 8) ·
Redissolution/remineralization
The rate of opal dissolution Ropal [mol Si m−3 yr−1 ] is proportional to the amount of opal in the sediment, copal ,
the pore water undersaturation with respect to silicic acid,
([Sisat ] − [Si]), and the rate constant, ropal :
Clim. Past, 7, 771–800, 2011
(A1)
· ([Sisat ] − [i]) · H ([Si]sat − [Si]).
Mopal stands for the molar weight of opal, ρ for the density,
8 for porosity and copal corresponds to the weight-fraction of
opal sediments. H (x) is the Heaviside step function, preventing spontaneous precipitation of solid opal in case of pore
water supersaturation. The silicic acid saturation concentration, [Si]sat , is a spatially constant model parameter set to
800 µ mol−l .
Since [Ca2+ ] is fairly constant and has a relatively high
background concentration, CaCO3 preservation is primarily governed by the concentration of carbonate ions in
the pore water. The rate of CaCO3 dissolution RCaCO3
[mol C m−3 yr−1 ] is thus linearly related to the pore water
undersaturation with respect to carbonate ions, ([CO3 ]sat −
[CO3 ]), the weight-fraction of CaCO3 , cCaCO3 , and the rate
constant, rCaCO3 :
ρ
MCaCO3
(A2)
× ([CO3 ]sat − [CO3 ]) · H ([CO3 ]sat − [CO3 ]),
2+ Ca
.
[CO3 ]sat =
0
Ksp
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T. Tschumi et al.: Deep ocean ventilation, carbon isotopes, marine sedimentation and the deglacial CO2 rise
0 is calculated after
The solubility product of calcite Ksp
Mucci (1983) with a pressure dependence following Millero
(1995). The calcium ion concentration [Ca2+ ] is proportional to salinity.
The rate of organic matter oxydation Roxy [mol C m−3
−1
yr ] is limited by the availability of oxygen in the pore water, [O2 ], and by the amount of POC, cPOC , in the sediment:
Roxy = roxy · cPOC · 8 · (1 − 8) ·
ρ
MPOC
·
RC
R O2
(A3)
· [O2 ] · H ([O2 ]).
The rate constant for oxydation, roxy , is set to 50 l mol−1
yr−1 . RC and RO2 are the Redfield ratios for carbon and
oxygen, respectively.
Organic matter denitrification only occurs in suboxic conditions, i.e. below a critical oxygen concentration in the pore
water [O2 ]crit . The reaction rate Rdenit [mol C m−3 yr−1 ] is
proportional to the POC weight-fraction, cPOC , the availability of nitrate, [NO−
3 ], and the rate constant rdenit :
Rdenit = rdenit · cPOC · 8 · (1 − 8) ·
ρ
RC
·
MPOC Rdenit
(A4)
Rdenit is the Redfield coefficient for denitrification. The
reaction rate, rdenit , is set to 50 l mol−1 yr−1 .
Inert clay is not subject to redissolution in the sediments:
Rclay = 0.
A2
cS corresponds to the dry weight fraction of CaCO3 , opal,
POC or clay. DB is the bioturbation coefficient, w is the
vertical advection velocity and Rx are the reaction rates.
The variable porosities are taken into account when solving
Eqs. (A6)–(A9).
Below the bioturbated zone, the sediment composition remains unchanged:
∂cS = 0, for each solid component S.
(A10)
∂z z≥10cm
At the sediment-water interface (z = 0 cm) the vertical advection velocity is given by the sum of the volume rain rates,
RainS · MS /ρ [cm3 cm−2 yr−1 ], divided by the solid sediment fraction, (1 − 8(0)):
X
1
w(0) =
RainS · MS /ρ.
(A11)
1 − 8(0) solid
components
Vertical advection velocities, w(z), decrease with depth
due to remineralization and redissolution as well as due to
increasing solid sediment fractions:
w(z) < w(0),
− · NO−
· H ([O2 ]crit − [O2 ]).
3 · H NO3
(A5)
−
The four solid sediment components are subject to bioturbation, advection and remineralization/redissolution. They
satisfy the following equations:
− RCaCO3 ·
MCaCO3
1
·
,
ρ
1−8
∂copal
∂2
∂
= DB 2 copal −
wcopal − Ropal
∂t
∂z
∂z
·
Mopal
1
·
,
ρ
1−8
∂2
∂cPOC
∂
= DB 2 (cPOC ) −
(wcPOC )
∂t
∂z
∂z
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X
solid
components

RS (z0 ) · MS (z0 )/ρ · dz0  .
The burial flux of the solid component S at the base of
the bioturbated zone (z = 10 cm), FSburial [mol cm−2 yr−1 ], is
given by:
ρ
· (1 − 8(10)).
MS
(A8)
(A9)
(A14)
In case total redissolution is larger than total rain in the
sediment column, gaps of solid material are produced within
the fixed model geometry. These are filled with clay in order
to ensure mass conservation:
copal + ccal + corg + cclay = 1.
MPOC
1
− Roxy + Rdenit ·
·
,
ρ
1−8
∂cclay
∂2
∂
= DB 2 cclay −
wcclay .
∂t
∂z
∂z
z0 =0

FSburial = w(10) · cS (10) ·
(A7)
(A12)
components
Zz
(A6)
for z > 0.
The velocity w(z) is given by the remaining volume flux
at depth z divided by the solid sediment fraction, (1 − 8(z)):

1
 X
w(z) =
RainS · MS /ρ
(A13)

1 − 8(z)
solid
Solid components
∂cCaCO3
∂2
∂
wcCaCO3
= DB 2 cCaCO3 −
∂t
∂z
∂z
795
A3
(A15)
Pore water solutes
The concentrations of pore water solutes are subject to sediment remineralization and redissolution as well as to pore
water diffusion. Here, the diffusion in a porous medium is
represented by multiplying the uniform pore water diffusion
coefficient D by the porosity 8. Experimentally derived diffusion coefficients do not deviate much from these values
Clim. Past, 7, 771–800, 2011
796
T. Tschumi et al.: Deep ocean ventilation, carbon isotopes, marine sedimentation and the deglacial CO2 rise
(Ullman and Aller, 1982). The dynamical equations
pore water solutes are:
∂Si
∂Si
∂
8D
+ Ropal /8,
=
∂t
∂z
∂z
∂O2
∂O2
R O2
∂
8D
−
=
· Roxy /8,
∂t
∂z
∂z
RC
!
RNO−
∂NO−
∂NO−
∂
3
3
3
=
8D
+
∂t
∂z
∂z
RC
for the
(A16)
(A17)
(A18)
Rdenit
· Rdenit /8,
RC
∂PO4
∂
∂PO4
RPO4
=
8D
+
(A19)
∂t
∂z
∂z
RC
· Roxy + Rdenit /8,
∂DIC
∂
∂DIC
(A20)
=
8D
∂t
∂z
∂z
+ Rcalc + Roxy + Rdenit /8,
RNO− + 2 · RPO4
∂ALK
∂
∂ALK
3
=
8D
−
(A21)
∂t
∂z
∂z
RC
· Roxy /8 −
· Roxy /8 +
Rdenit − 2 · RPO4
· Rdenit /8
RC
+ 2 · Rcal /8.
The boundary condition for pore water solutes at the
ocean-sediment interface is:
C|z=0cm = Cbottom water ,
for each solute tracer C. (A22)
Below the bioturbated zone the solute concentrations are
assumed to remain constant:
∂C = 0, for each solute tracer C.
(A23)
∂z z≥10cm
The pH-equilibrium in the pore water is given by:
[CO2 ] CO=
K0
3
= 10 .
(A24)
2
K2
HCO−
3
The apparent dissociation constants K10 and K20 are calculated after Mehrbach et al. (1973) with pressure dependence
following Millero (1995).
Acknowledgements. This study received support by the Swiss
National Science Foundation, by the Swiss Staatssekretariat für
Bildung und Forschung (#C07.0068; COST Action 735), and by
the European Commission through the FP7 projects Past4Future
(grant no. 243908), CARBOCHANGE (grant no. 264879), and
EPOCA (grant no. 211384). This is publication no. A 343 of
the Bjerknes Centre for Climate Research. The comments by
T. F. Stocker are greatly appreciated.
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